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7. Mountain Glaciers and Ice Caps

Coordinating Lead Author: Martin Sharp

Lead Authors: Maria Ananicheva, Anthony Arendt, Jon-Ove Hagen, Regine Hock, Edward Josberger, R. Dan Moore, William Tad Pfeffer, Gabriel J. Wolken

Contributing Authors: Helgi Björnsson, Carl Egede Bøggild, Tobias Bolch, Igor Buzin, John J. Clague, John Graham Cogley, Julian A. Dowdeswell, Mark B. Dyurgerov, Andrey Glazovsky, Hester Jiskoot, Tómas Jóhannesson, Alexander Klepikov, Alexander Krenke, Mark F. Meier, Brian Menounos, Alexander Milner, Yaroslav Muravyev, Matt Nolan, Finnur Palsson, Valentina Radi ´c, Mattias de Woul

Contents

Key Findings . . . . 2

Summary . . . . 2

7.1. Introduction . . . . 3

7.1.1. Background . . . . 3

7.1.2. Context: What did the Arctic Climate Impact Assessment report say about mountain glaciers and ice caps? . . . . 4

7.1.3. Geographic setting . . . . 4

7.1.4. Characteristics of Arctic mountain glaciers and ice caps . . 5

7.1.5. Significance and impacts of glacier changes . . . . 7

7.1.6. Challenges . . . . 8

7.2. Climate evolution in glacierized regions of the Arctic . . 8

7.2.1. Holocene climate . . . . 8

7.2.2. Arctic ice cores: Holocene climate records from glacierized regions . . . . 9

7.2.3. Climate records for glacierized regions of the Arctic derived from instrumental observations and climate re-analysis . . . . 10

7.3. Changes in glacier extent, volume and total mass . . . . 11

7.3.1. Arctic glacier changes during the Holocene . . . . 11

7.3.2. Arctic glacier changes in the 20th and 21st centuries . . 12

7.3.3. Links between mass balance and climate . . . . 18

7.4. Proxy indicators of surface mass balance . . . . 18

7.4.1. Introduction . . . . 19

7.4.2. Proxy indicators, data sources, and methods of measurement . . . . 19

7.5. Ice dynamics and iceberg calving . . . . 22

7.5.1. Overview . . . . 23

7.5.2. Measurement methods . . . . 24

7.5.3. Calving and surging glaciers in the Arctic . . . . 25

7.5.4. Recent ice shelf break-up events . . . . 27

7.5.5. Iceberg characteristics and relationship to ice dynamics . 28 7.5.6. Controls on calving fluxes . . . . 28

7.5.7. State of theory and modeling . . . . 28

7.5.8. Requirements for improved predictions of calving fluxes . 29 7.6. Projections of Arctic glacier changes . . . . 29

7.6.1. Downscaling climate model projections . . . . 29

7.6.2. Modeling mountain glacier and ice cap mass balance in the 21st century . . . . 30

7.6.3. Modeling ice dynamics and ice extent . . . . 31

7.7. Impacts of changes in mountain glaciers and ice caps 33 7.7.1. Impacts on sea level . . . . 33

7.7.2. Impacts on the marine environment . . . . 34

7.7.3. Impacts on water resources . . . . 36

7.7.4. Glacier ecosystems . . . . 39

7.7.5. Geomorphological hazards . . . . 40

7.7.6. Glacier-related tourism . . . . 41

7.8. New information expected from International Polar Year projects . . . . 42

7.8.1. Alaska . . . . 42

7.8.2. Arctic Canada . . . . 42

7.8.3. Iceland . . . . 43

7.8.4. Svalbard . . . . 43

7.8.5. Russian Arctic and mountains . . . . 43

7.9. Synthesis of the current state of Arctic mountain glaciers and ice caps . . . . 44

7.9.1. Mountain glacier and ice cap mass balance . . . . 44

7.9.2. Ablation due to iceberg calving . . . . 45

7.9.3. Observed trends in ice extent . . . . 46

7.9.4. Relationship to climate and circulation changes . . . . 46

7.9.5. Projections of future change . . . . 47

7.9.6. Impacts . . . . 47

7.10. Knowledge gaps and recommendations . . . . 47

7.10.1. Key gaps in knowledge . . . . 48

7.10.2. Basic observations . . . . 48

7.10.3. Mass balance measurements, proxies, and modeling . . 48

7.10.4. Ice dynamics and ablation by iceberg calving . . . . 49

7.10.5. Impacts . . . . 49

Appendix 7.1 Glossary of glaciological terms. . . . 50

References . . . . 52

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Key Findings

• Mountain glaciers and ice caps in the Arctic cover an area of around 402 000 km2 and contain about 0.41 m sea-level equivalent of water.

• Over the past century, nearly all have retreated from maximum extents reached during the Little Ice Age, which ended in the late 19th century. This period of glacier retreat has been associated with an overall reduction in glacier mass during the period of record, which extends to more than 60 years in some cases.

Surface mass balance measurements showed negative or nearly balanced conditions and generally showed no trend until the mid-1990s, but since then reveal significantly higher rates of mass loss in Alaska, the Canadian Arctic, and Iceland.

• The fraction of ablation that occurs through iceberg calving can be as much as 40% in regions where it has been measured, but it has not been measured over large areas of the Arctic. Estimation of calving fluxes is therefore a major source of uncertainty in estimates of current and future rates of mass loss from mountain glaciers and ice caps in the Arctic.

• Mass loss (surface mass balance plus calving) from Arctic glaciers probably exceeded 150 Gt/y in the past decade, when it was similar to mass loss from the Greenland Ice Sheet. This suggests that glacier and ice sheet change in the Arctic is probably now the dominant contributor to the eustatic (water mass) component of global sea level rise.

• Under the IPCC A1B emissions scenario, the total volume of Arctic glaciers is projected to decline by between 13% and 36% by 2100 (corresponding to an increase of 51 to 136 mm sea-level equivalent), depending on the choice of general circulation model.

These projections are a lower bound since they do not include mass losses by iceberg calving. Regardless, mountain glaciers and ice caps will continue to influence global sea-level changes beyond the 21st century.

• In many parts of the Arctic, climate warming should cause glacier runoff to increase for a few decades or longer, but glacier area reduction will ultimately cause glacier runoff to decline. These changes in glacier runoff will have impacts on water supplies; water quality; hydroelectric power generation; flood hazards;

freshwater, estuarine and coastal habitats; and ocean circulation patterns.

• Iceberg hazards to shipping and offshore activities related to exploration for and exploitation of offshore hydrocarbon and mineral resources may increase if changes in tidewater glacier dynamics result in more iceberg production and/or larger bergs, and reductions in sea-ice cover allow icebergs to become more mobile.

Summary

In addition to the Greenland Ice Sheet, the Arctic contains a diverse array of smaller glaciers ranging from small cirque glaciers to large ice caps with areas up to 20 000 km2. Together, these glaciers cover an area of more than 400 000 km2, over half the global area of mountain glaciers and ice caps. Their total volume is sufficient to raise global sea level by an average of about 0.41 m if they were to melt completely.

These glaciers exist in a range of different climatic regimes, from the maritime environments of southern Alaska, Iceland, western Scandinavia, and Svalbard, to the polar desert of the Canadian Arctic. Glaciers in all regions of the Arctic have decreased in area and mass as a result of the warming that has occurred since the 1920s (in two pulses – from the 1920s to the 1940s and since the mid-1980s). A new phase of accelerated mass loss began in the mid-1990s, and has been most marked in Alaska, the Canadian Arctic, and probably Greenland. Current rates of mass loss are estimated to be in the range 150 to 300 Gt/y; comparable to current mass loss rates from the Greenland Ice Sheet. This implies that the Arctic is now the largest regional source of glacier contributions to global sea-level rise.

Most of the current mass loss is probably attributable to a change in surface mass balance (the balance between annual mass addition, primarily by snowfall, and annual mass loss by surface melting and meltwater runoff). Iceberg calving is also a significant source of mass loss in areas such as coastal Alaska, Arctic Canada, Svalbard, and the Russian Arctic. However, neither the current rate of calving loss nor its temporal variability have been well quantified in many regions, so this is a significant source of uncertainty in estimates of the total rate of mass loss. It is, however, clear that the larger Arctic ice caps have similar variability in ice dynamics to that of the Greenland Ice Sheet. That is to say, areas of relatively slow glacier flow (which terminate mainly on land) are separated by faster-flowing outlet glaciers (which terminate mainly in the ocean). Several of these outlet glaciers exhibit surge-type behavior, while others have exhibited substantial velocity changes on seasonal and longer timescales. It is very likely that these changes in ice dynamics affect the rate of mass loss by calving both from individual glaciers and the total ice cover.

Projections of future rates of mass loss from mountain glaciers and ice caps in the Arctic focus primarily on projections of changes in the surface mass balance. Current models are not yet capable of making realistic forecasts of changes in losses by calving. Surface mass balance models are forced with downscaled output from climate models driven by forcing scenarios that make assumptions about the future rate of growth of atmospheric greenhouse gas concentrations. Thus, mass loss projections vary considerably, depending on the forcing scenario used and the climate model from which climate projections are derived. A new study in which a surface mass balance model is driven by output from ten general circulation models (GCMs) forced by the IPCC (Intergovernmental Panel on Climate Change) A1B emissions scenario yields estimates of total mass loss of between 51 and 136 mm sea-level equivalent (SLE) (or 13% to 36% of current glacier volume) by 2100. This implies that there will still be substantial glacier mass in the Arctic in 2100 and that Arctic mountain glaciers and ice caps will continue to influence global sea-level change well into the 22nd century.

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As glaciers and ice caps shrink in a warming climate, runoff initially increases in response to higher rates of surface melting. Ultimately, however, runoff will decline as reductions in glacier area outweigh the effect of more rapid melting.

This phase of declining runoff does not yet seem to have begun in most regions of the Arctic, but it may begin soon in the Russian Arctic mountains. In the Yukon River basin, Greenland, Iceland, and Norway, glacier runoff is an important resource for hydroelectric power generation, and the viability of hydroelectric projects may ultimately be compromised by runoff decreases associated with glacier shrinkage.

Changes in glacier runoff also result in changes in stream temperature, sediment load, and nutrient export (both magnitude and type) that can be expected to initiate changes in the ecology and productivity of downstream river, lake, and fjord environments. Increased rates of glacier melt may accelerate the release of a range of ‘legacy’ pollutants stored in firn (partially compacted snow that is the intermediate stage between snow and glacier ice) and glacier ice back into the environment. Increasing freshwater fluxes to fjords and other nearshore marine environments will alter the characteristics of surface water masses and drive changes in circulation.

Circulation changes may also follow the retreat of tidewater glaciers onto land, particularly in regions of upwelling close to the termini of these glaciers. These changes can decrease the availability of feeding and resting habitats that are important for marine mammals and seabirds.

The number and size of icebergs produced is likely to change as tidewater glaciers retreat, ultimately reaching zero as their termini emerge onto land. Break-up of floating glacier tongues and ice shelves, a process that has accelerated in the past decade along the northern coast of Ellesmere Island, results in large tabular bergs, while accelerated flow of tidewater glaciers tends to result in accelerated production of small bergs unless flotation of the glacier terminus occurs, when large tabular bergs may be produced. Cessation of small berg production when tidewater termini retreat onto land can reduce the number of such bergs that become grounded in fjords, decreasing the availability of important resting habitat for seals. Circulating icebergs are a potential hazard for shipping, drilling platforms, and seafloor pipelines in the Arctic. Circulation patterns and longevity of bergs may change as the Arctic sea-ice cover declines. Thinner and less extensive sea ice is likely to result in greater berg mobility, while warmer surface waters may result in more rapid melting and disintegration of bergs. A knowledge of the size distributions of bergs produced and circulating in different regions is critical to evaluating the risk of bergs contacting the sea floor and damaging seafloor pipelines.

Glacier retreat will be associated with changes in the magnitude and frequency of a range of geomorphological hazards, most notably outburst floods from ice-marginal and moraine-dammed proglacial lakes, and mass movements from newly deglaciated valley walls. The degree of risk from such phenomena will, however, be highly variable, depending upon the nature of the deglaciated terrain, the size of local populations, and the amount of infrastructure present in individual regions.

Finally, it should be emphasized that the ability to monitor and predict changes in the Arctic’s mountain glaciers and ice caps is still quite limited. Basic inventory data for Arctic

glaciers are lacking. The number of glaciers on which mass balance is measured is small and declining, and the distribution of measurement sites is highly non-uniform. There are no measurement sites at all in some areas with large glacier areas (such as the Russian Arctic and the Yukon). There is no routine monitoring of mass losses by iceberg calving, and understanding of what controls calving rates is rudimentary.

This severely constrains the ability to model calving losses into the future. Studies of the socio-economic impacts of Arctic glacier change are currently few and limited in scope, so most statements about such impacts are based solely on general principles. As such, they do not provide a strong basis for either the formulation or enactment of a policy response to Arctic glacier change. Some of the largest impacts of the ongoing changes in Arctic glaciers (such as global sea-level rise) will be felt in regions of the world that are very far from the Arctic.

They may, however, still have social, political, and economic repercussions for Arctic nations – repercussions that need to be explored more thoroughly.

7.1.

Introduction

• Mountain glaciers and ice caps cover an area of nearly 402 000 km2 in the Arctic. Over half of this area is in western North America and the Canadian Arctic.

• The combined volume of these glaciers is sufficient to raise sea level by around 0.41 m if they all melted.

• Glacier types range from small cirque glaciers to large ice caps, with areas of up to 20 000 km2. These ice caps are dynamically complex, and are drained in part by fast- flowing outlet glaciers that often reach the ocean and lose mass by calving icebergs.

• In many regions, a proportion of the glaciers exhibit ‘surge- type’ behavior, in which long periods of flow at relatively low speeds are punctuated by short-lived episodes of very rapid flow.

• Long-term changes in the extent, volume, and mass of these glaciers are driven by changes in climate and oceanographic conditions that alter their ‘mass balance’ – the annual balance between mass gains (due mainly to snowfall) and mass losses (due mainly to surface melting and runoff, and iceberg calving).

• Changes in the thermal structure and flow of glaciers can play an important role in how, and how rapidly, they respond to changing climate and oceanographic conditions.

7.1.1.

Background

Glacier ice (including mountain glaciers, ice caps and the Greenland Ice Sheet) occupies about 2.16 million km2 of the Arctic land surface. This chapter deals with all mountain glaciers and ice caps (hereafter referred to as glaciers), including those in Greenland that are not connected to the ice sheet. The Greenland Ice Sheet itself is addressed in Chapter 8. Mountain glaciers are ice bodies whose geometry and boundaries are controlled by bedrock topography, while ice caps are dome- shaped ice bodies that submerge the underlying bedrock topography. Table 7.1 presents the regional distribution of the

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area and volume of glacier ice in the Arctic. Here, the ice volume is given in units of mm sea-level equivalent (SLE), which is the volume of water (in m3) stored as glacier ice, divided by the surface area of the global ocean (in m2) multiplied by 1000.

7.1.2.

Context: What did the Arctic Climate Impact Assessment report say about mountain glaciers and ice caps?

The SWIPA report is an update of knowledge on the state of the Arctic cryosphere relative to the Arctic Climate Impact Assessment (ACIA, 2005). This dealt with mountain glaciers, ice caps, and the Greenland Ice Sheet in a single section of chapter 6, ‘Cryosphere and Hydrology’. It gave a regional summary of glacier changes since around 1950 with an emphasis on measurements of surface mass balance and, for regions where mass balance measurements were lacking, changes in glacier area. Mass losses by iceberg calving were mentioned but not discussed in detail. The major ACIA contribution was a series of projections of changes in the surface mass balance of Arctic glaciers for the period until 2100. These mass balance projections were computed using regional seasonal sensitivity characteristics (a measure of the expected change in the annual surface mass balance due to prescribed changes in monthly mean air temperature and precipitation amounts) and projected changes in temperature and precipitation derived from a suite of five GCMs forced by the IPCC B2 emissions scenario. The mass balance projections were used to estimate the potential contributions of different regions of the Arctic to sea level rise over the 21st century.

The ACIA report also commented on the potential impacts of changes in glaciers for other parts of the physical system,

ecosystems, and people, and gave an assessment of critical research needs that remains valid today.

7.1.3.

Geographic setting

The glacierized areas (i.e. those presently overlain by glaciers) addressed in this chapter are shown in Figure 7.1. Most of the glaciers are located north of 60° N, but some more southerly glaciers, in Kamchatka, Alaska, and northwestern Canada are also included (Table 7.1). The total glacier-covered area is approximately 402 000 km2, which is about 54% of all glaciers in the world excluding the large ice sheets (~741 400 km2, Radi´c and Hock, 2010).

Arctic glaciers are irregularly distributed in space (Figure 7.1), and are located in a range of very different climatic regimes (Braithwaite, 2005). In southern Alaska, Iceland, western Scandinavia, and western Kamchatka, the climate is maritime with a relatively small annual temperature range and precipitation rates of a few metres per year, while in the Canadian High Arctic it is very dry, cold and continental, with short summers, a very large annual temperature range (greater than 50 °C), and annual precipitation that ranges from 0.1 to 0.7 m/y. Conditions on Svalbard and in the Russian Arctic islands fall between these two climatic extremes. In northwestern North America, most of the ice cover is in the mountain ranges adjacent to the Gulf of Alaska, while in Arctic Canada the most heavily glacierized regions are in the mountains on Devon and Ellesmere Islands, which are nourished in part by moisture from a large persistent polynya in northern Baffin Bay (the North Open Water).

Region Ice-covered area, km2 Volumea, mm SLE Source, area

Canadian Arctic 1 51 000 199 ± 30 Ommanney, 1970

Northwestern North America 91 800 71 ± 8 Berthier et al., 2010; ESRI, 2003

Russian Islandsb 56 700 44 ± 8 Radi´c and Hock, 2010

Greenlandc 48 600 44 ± 7 Weng, 1995

Svalbard 36 500 26 ± 2 Hagen et al., 1993

Iceland 11 000 12 ± 6 Björnsson and Pálsson, 2008

Scandinavia 3 100 6 ± 0 Østrem et al., 1973, 1988

Russian Arctic 2 900 4 ± 0 Radi´c and Hock, 2010

Total (Arctic) 401 600 410 ± 30

Mountain glaciers and ice caps (global) 741 400 600 ± 70 Radi´c and Hock, 2010

Greenland Ice Sheet 1 755 600 7500 Bamber et al., 2001

Antarctic ice sheet

Antarctic ice shelves 12 348 000

1 555 000 57 000 Fox and Cooper, 1994 (areas)

Lythe et al., 2001 (volume)

Table 7.1. The areas (rounded to three significant figures) and estimated volumes, in mm sea-level equivalent (SLE), of ice caps and glaciers in the Arctic compared to global values and the ice sheets in Greenland and Antarctica.

a Volumes of all mountain glaciers and ice caps are from Radi´c and Hock (2010); b Franz Josef Land (13 700 km2), Novaya Zemlya (23 600 km2), Severnaya Zemlya (19 400 km2); c Greenland excluding the ice sheet: only the glaciers physically disconnected from the ice sheet are considered here.

Areas over 70 000 km2 have been reported (Holtzscherer and Bauer, 1954; Weidick and Morris, 1998); however, these estimates have included some glaciers that are connected to the ice sheet but considered independent of the ice sheet in a dynamic sense or that were, by their morphology, discernible as units independent of the ice sheet. Nevertheless, the estimate by Weng (1995) is a minimum estimate because it is based on a 1:2 500 000 map, a scale too coarse to allow identification of many small glaciers.

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7.1.4.

Characteristics of Arctic mountain glaciers and ice caps

7.1.4.1.

Morphology

Arctic glaciers have a range of forms. Dome-shaped ice caps have lobes and outlet glaciers that drain ice away from the accumulation area, where annual snowfall exceeds annual surface melt, to lower-lying regions (ablation areas) where the reverse is true, and in some cases to tidewater margins where the ice reaches the ocean and icebergs are calved. Large ice caps are found in the Canadian Arctic, Iceland, Svalbard, the Russian Arctic islands, and in Greenland beyond the margins of the ice sheet (Figure 7.2). In other regions, large glaciers originate from icefi elds that fi ll basins within mountain ranges (e.g., in southern Alaska). Many regions, for example Svalbard, also have a large number of individual valley glaciers that occupy valleys and basins in the landscape.

Much of the Arctic ice mass is contained in relatively large ice caps with areas of up to 20 000 km2, although there are large numbers of independent glaciers with areas ranging from 0.1 to several thousand km2 (Dowdeswell and Hagen, 2004). The largest ice masses in the Arctic are found in Arctic Canada, and include the Agassiz Ice Cap (17 300 km2) and Prince of Wales Icefi eld (19 400 km2) on Ellesmere Island, the Devon Island Ice Cap (about 14 400 km2), and the Barnes and Penny Ice Caps (each almost 6000 km2) on Baffi n Island. On Greenland, the largest independent ice cap is the ~9000 km2 Flade Isblink in the northeast, while the more northerly Hans Tausen Ice Cap,

at 3975 km2, is probably the best studied (Reeh et al., 2001).

Austfonna on Nordaustlandet in eastern Svalbard (8120 km2) is the largest ice cap in the Eurasian Arctic (Dowdeswell, 1986;

Hagen et al., 1993). The largest Russian ice cap is the Academy of Sciences Ice Cap (5575 km2) on Severnaya Zemlya (Dowdeswell et al., 2002), although the northern island of Novaya Zemlya has a larger ice-covered area with many outlet glaciers. In Iceland, Vatnajökull has an area of 8100 km2 (Björnsson and Pálsson, Scandinavia

Svalbard

Iceland Greenland

Arctic Canada Northwestern North America

Russian islands

Russian Arctic

Pan-Arctic glaciers Pan-Arctic glacier regions Greenland Ice Sheet (considered in Chapter 8) Arctic Circle

Figure 7.1. Map of the Arctic showing the distribution of regions containing mountain glaciers and ice caps considered in this report and the major glacierized areas.

Figure 7.2. Small plateau ice caps on Axel Heiberg Island, Arctic Canada, with a larger ice cap in the distance. Note the clear contrast between the snow-covered accumulation area and the greyer ablation area, and the outlet glaciers draining into valleys leading away from the ice cap. Source:

Martin Sharp, University of Alberta.

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2008). Among the mountain glaciers (Figure 7.3), the largest is Bering Glacier in Alaska, with an area of 3630 km2.

The larger Arctic ice caps and icefields (Figure 7.4) have complex dynamics involving a mix of fast- and slow-flowing elements that can vary in how they respond to climate changes and variability. Vestfonna on Nordaustlandet in eastern Svalbard is a good example, where four fast-flowing outlet glaciers are embedded within a largely slower-flowing ice cap of 2500 km2 (Dowdeswell and Collin, 1990). Typical surface velocities may be less than 10 m/y close to the equilibrium-line altitude of valley and cirque glaciers that terminate on land. They can reach many hundreds of metres per year on major calving outlet glaciers, on which seasonal velocity variations can be large, with summer velocities up to an order of magnitude greater than winter velocities (Williamson et al., 2008). Interannual velocity variations can also be significant on such glaciers.

7.1.4.2.

Thermal regime

Most glaciers in Iceland, Scandinavia, and central and southern Alaska are predominantly temperate (composed of ice at the pressure melting point). Elsewhere in the Arctic, glaciers tend

to have polythermal temperature regimes (composed of a mixture of ice at and below the pressure melting point), and their dynamics may be strongly affected by climate-driven changes in the thermal regime. In some parts of polythermal glaciers, the ice temperature at the glacier bed is below the melting point, implying that the glacier is frozen to its bed.

In some smaller (cold) glaciers, all the ice is at temperatures below the melting point, except at the surface where the ice temperature may reach the melting point in summer. The thermal regime of glaciers is determined by the prevailing climatic conditions (snow accumulation and surface energy balance). Climatic change can alter the thermal regime of a glacier, and potentially also its dynamics, because ice deforms more rapidly at higher temperatures and glacier flow can be enhanced by the lubricating effect of meltwater at the glacier bed. Most cold-based and polythermal glaciers are found in dry regions with low accumulation rates. It takes a long time for a climate change signal to penetrate into such glaciers, and changes in the temperature regime are probably not very large on a 100-year timescale. However, in areas where penetration of surface meltwater into cold firn on the glacier surface increases, the release of latent heat when this meltwater refreezes can cause a more rapid change in the thermal regime.

Owing to the polythermal nature of many Arctic glaciers, the formation of superimposed ice by meltwater refreezing on the glacier surface and internal accumulation (where percolating meltwater freezes in cold snow and firn) can be important processes of mass accumulation. When these processes occur, the refrozen water has to be melted again to become meltwater runoff, complicating the measurement and modeling of the surface mass balance. Neglecting or inadequately accounting for these processes in mass balance measurements overestimates mass loss. Superimposed ice formation is important on many Arctic glaciers and is the dominant form of accumulation on some (Koerner, 1970).

Surge-type glacier behavior is common in many parts of the Arctic and can be a source of significant hazards. In a glacier surge, surface velocities increase by an order of magnitude or more and glacier fronts can advance many kilometres (sometimes more than 10 km) in a matter of years. When tidewater glaciers surge, iceberg production increases and can be an important process of mass loss from the glacier (Liestøl, 1973; Dowdeswell, 1989). While the occurrence of individual surges is not directly related to climate change, climate change may alter the frequency of surging (Dowdeswell et al., 1991;

Eisen et al., 2001), or even cause glaciers to cease being of the surge type (Dowdeswell et al., 1995). Changes in the number of actively surging glaciers in a region that extend into tidewater can have a large short-term impact on the regional glacier mass balance through changes in the overall calving flux to the ocean.

7.1.4.3.

Tidewater glaciers

Iceberg calving plays a major role in the overall mass balance of tidewater glaciers and Arctic ice caps that have significant tidewater margins (Dowdeswell et al., 1997, 2008; Dowdeswell and Hagen, 2004; Błaszczyk et al., 2009) (Figure 7.5). However, accurate calculation of the calving flux requires information that is frequently not available (e.g., ice thickness at the glacier terminus), and modeling of calving flux is challenging because Figure 7.4. The Glacier Bay Icefield, Alaska. Outlet glaciers drain outward

from a central accumulation area within the mountains. Source: Landsat imagery courtesy of NASA Goddard Space Flight Center and U.S.

Geological Survey.

Figure 7.3. Gulkana Glacier in Alaska is a typical mountain glacier and a long-term mass balance monitoring site. Note the end moraines surrounding the glacier terminus that indicate that the glacier was more extensive and thicker in the past. Source: U.S. Geological Survey.

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of the lack of a widely accepted calving law. Although individual ice caps may have multiple tidewater glacier outlets, the mass loss by calving is often dominated by one or two of these outlets (Krenke, 1982; Dowdeswell et al., 2002, 2008; Burgess et al., 2005; Williamson et al., 2008; Mair et al., 2009), but calving fluxes from individual glaciers can show large temporal variability.

Recent increases in calving flux from the Greenland Ice Sheet have been associated with acceleration, retreat, and thinning of major outlet glaciers (Howat et al., 2005; Rignot and Kanagaratnam, 2006). Similar changes are observed in Alaska and on some large Arctic ice caps, and may be caused by changes in glacier dynamics related to increased surface melting, penetration of meltwater to glacier beds, and subsequent lubrication of the ice-bed interface allowing increased flow by basal sliding. Other possible causes of calving flux increases include increased melt of the underwater part of terminal ice cliffs due to increasing ocean temperature (Holland et al., 2008), thermal transitions (from cold-based to warm- based) in polythermal glaciers, and break-up and removal of floating ice tongues that formerly restrained the flow of the glacier. Such dynamic changes may be larger and more rapid than those induced by changes in surface mass balance alone.

7.1.4.4.

Response to climate change

Glaciers respond to climate change over very different timescales depending on their size, shape, and thermal regime.

Among glaciers that terminate on land, smaller glaciers tend to respond more quickly, changing their shape, flow, and terminus position over years or decades. The hypsometry (area-altitude distribution) of a glacier plays an important controlling role in determining its response to changes in climate. In a given region, low-lying glaciers may shrink and retreat quickly at the same time as glaciers at higher elevations grow or maintain their size. Changes in the temperature and salinity of the ocean water adjacent to calving ice fronts can trigger rapid changes in the terminus position, flow velocity, and calving flux of tidewater glaciers, so that large tidewater glaciers may alter their form more rapidly than small glaciers that terminate on

land (Holland et al., 2008; Rignot et al., 2010; Straneo et al., 2010). Although Arctic glaciers show a variety of responses to changing climate, it is something to which they are all sensitive.

It is worth noting that the surface mass balance of glaciers in the Arctic may be affected indirectly by changes in other components of the Arctic cryosphere, such as the regional snow cover and extent of sea ice, that affect the surface energy balance, atmospheric circulation, and availability of open water as a source of water vapor (e.g., Rennermalm et al., 2009).

7.1.5.

Significance and impacts of glacier changes

According to recent estimates, glaciers contributed about 0.8 to 1 mm annually to global sea-level rise during the period 2001 to 2004. The upper limit of the estimates includes the contribution from glaciers surrounding the Greenland and Antarctic ice sheets (Kaser et al., 2006; Solomon et al., 2007). Meier et al. (2007) gave a slightly higher estimate of 1.1 mm for 2006. This is about one third of the total current sea-level rise, or about 60% of that component of the rise attributable to the addition of water mass to the oceans, as opposed to the thermal expansion of ocean waters. The remaining 40% of that component comes from the combined contribution of the Greenland and Antarctic ice sheets. Meier et al. (2007) predicted that a large contribution to sea-level rise over the period to 2100 will still come from glaciers and ice caps. As more than half of the world’s glacier and ice cap area is found in the Arctic, it is very important to reduce the uncertainty in estimates of the current and future mass balance of the Arctic glaciers and ice caps.

It is also important to evaluate the impacts of ongoing and future changes in glacier extent and volume on regional water resources, water quality, and the incidence of glacier-related natural hazards. Glacier runoff is exploited as a source of hydroelectric power in Scandinavia, Iceland, Greenland, western Canada, and Alaska. Increasingly negative surface mass balance of the glaciers, and reductions in glacier area may have a direct impact on the water balance of basins used for hydroelectric power generation. Increased freshwater flux from glaciers to nearby fjords and oceans may have an impact on the marine ecosystem via freshening of ocean water and increased transport of nutrients and contaminants from land to sea. Freshening of ocean water may have an impact on fjord circulation and on global ocean thermohaline circulation. Changes in the frequency and magnitude of iceberg calving events may impact infrastructure development, marine transport, fisheries, and oil and gas exploration and production on Arctic continental shelves, where human activity is expected to increase significantly over the coming decades.

Changes in the extent and thickness of Arctic glaciers can destabilize the surrounding terrain and generate geomorphological hazards (rock slides, debris flows, ice avalanches, glacial mudflows, outburst floods), especially when combined with changes in the extent and thickness of permafrost on surrounding slopes. They also create a threat of floods (jökulhlaups) from ice-dammed and subglacial lakes, especially in volcanic areas such as Iceland, Alaska, and Kamchatka.

Glacier related tourism is important to some local economies, and shrinking or even disappearing glaciers may have a direct economic impact.

Figure 7.5. Time-lapse camera overlooking the calving terminus of the tidewater glacier, Kronebreen, in Svalbard. Kronebreen is a focus of study within the IPY GLACIODYN project. Source: Monica Sund, University Centre, Svalbard.

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7.1.6.

Challenges

The ACIA report (ACIA, 2005) stated that the most difficult task in a regional-scale assessment of glacier behavior is to generalize results from a few glaciers and ice caps to all ice masses in the Arctic. This statement is still valid today. Neither the remote sensing tools nor the glacier and ice sheet models currently available are well suited to studies of regional ice covers that comprise a multitude of glaciers of varying sizes in regions with complex surface topography.

Direct measurements of the mass balance of Arctic glaciers are limited to a small number of glaciers across the Arctic (50 in the 2000 to 2004 pentad); 31 of these were in Scandinavia, which contains only 3.5% of the total mountain glacier and ice cap area in the Arctic. In the past 35 years, the number of in situ glacier monitoring sites in the Canadian Arctic and Arctic Russia (regions which contain over 50% of the area of mountain glaciers and ice caps in the Arctic) has declined from 16 in the 1970 to 1975 period, to just six in the period 2000 to 2005, with no measurements in Arctic Russia since 1990 and only one set since 1980 (Dyurgerov and Meier, 2005). There is a pressing need for an updated regional-scale assessment of glacier and ice cap mass balance, the last assessment having been made in 2005. The uncertainties associated with such regional-scale assessments are large owing to the small number of in situ measurements and their uneven spatial distribution relative to the distribution of glacier ice.

In situ measurements are critical for quantifying regional mass balance, enhancing process understanding, and validating remote sensing techniques and model predictions. Given the limited number and distribution of such measurements, it is essential to improve and sustain remote sensing capabilities for monitoring ongoing changes in glacier extent, surface elevation and thickness, surface mass balance, ice dynamics, and iceberg production. Although there has been some progress on these issues, there is also a need for repeated, regional-scale mapping of parameters that provide simple indices of glacier mass balance (such as equilibrium-line altitude, glacier facies zones, and summer melt duration; see Section 7.4.2) to be used to evaluate year-to-year variability in climate effects on glacier health.

7.2.

Climate evolution in glacierized regions of the Arctic

• While some Arctic glaciers have existed throughout the period since the end of the last glaciation, most disappeared during a warm period between 10 000 and 6000 years ago. Many of today’s glaciers, therefore, formed in a cooler period after 5000 years ago.

• Although there were several warmer intervals within this cool period, the 20th century appears to have been the warmest century in the past 2000 years, with the warmest conditions occurring between the 1930s and early 1960s and since the mid-1980s.

• Across much of the Arctic, low winter precipitation means that most of the year-to-year variability in glacier mass balance arises from changes in summer temperature. In more maritime regions such as southern Alaska, Iceland, western Scandinavia and Svalbard, variability in winter precipitation can be an important influence on mass balance variability.

7.2.1.

Holocene climate

Ice formed during the last glaciation is found in some ice caps in the Canadian Arctic (Koerner and Fisher, 2002; Zdanowicz et al., 2002) and Severnaya Zemlya (Kotlyakov et al., 1991). The absence of such ice in ice caps in Svalbard and other parts of the Russian Arctic is evidence for substantial retreat (or even disappearance) of glaciers and ice caps across much of the Arctic outside Greenland in the early Holocene (Koerner and Fisher, 2002). Therefore, the genesis of much of the present- day Arctic glacier cover may lie in the climate of the middle and late Holocene.

Many climate proxy records suggest that the warmest period of the Holocene in the Arctic was between 10 000 and 6000 years bp, after which the climate cooled, possibly because of a decrease in incident solar radiation at high latitudes due to changes in the Earth’s orbital parameters (Koerner and Fisher, 1990; Vinther et al., 2009). Air temperature reconstructions based on ice-core oxygen isotope records from the Agassiz Ice Cap (Canadian Arctic) and Renland (eastern Greenland) suggest peak temperatures slightly in excess of 2 °C warmer than at present during this period (Vinther et al., 2009), which is often referred to as the Holocene Thermal Maximum. However, the exact timing of peak air temperatures during this period seems to have varied across the Arctic. In North America, it was later in the eastern Arctic than in the western Arctic, probably because the residual Laurentide Ice Sheet had a cooling effect on the climate in the east. During the Thermal Maximum, around 7500 years bp, summer temperatures were about 1.6 ± 0.8 °C warmer than in the 20th century (Kaufman et al., 2004).

Records of summer melt and oxygen isotope ratios from ice cores in the Canadian and Russian Arctic indicate that cooling following the Thermal Maximum was underway by no later than 6000 years bp (Koerner and Fisher, 2002) (Figure 7.6).

A 2000-year, decadally-resolved, multi-proxy record of Arctic summer surface temperatures (which includes the temperature reconstructions from Agassiz Ice Cap and Renland) shows a cooling trend from 2000 to 100 years bp (Kaufman et al., 2009) (see Fisher, 2002, for an analysis of the reliability of multi-proxy temperature reconstructions).

This trend (-0.22 °C/millennium) was probably driven by the orbitally-forced trend in summer insolation at high latitudes over the same period, but may have been amplified in some regions by feedbacks involving the progressive expansion of Arctic sea ice (England et al., 2008; Vare et al., 2009) and seasonal snow cover. The record also shows centennial-scale anomalies about the long-term trend, with the period 450 to 700 ad being cooler than the long-term trend and the period 900 to 1050 being warmer. The coldest period in the record occurred between 1600 and 1850, during the Little Ice Age (Kaufman et al., 2009). There is evidence that this cold period was associated with a persistently negative phase of the North Atlantic Oscillation (Trouet et al., 2009). Although the decline in orbitally-driven summer insolation continued throughout the 20th century, reconstructed Arctic summer temperatures rose sharply, reaching the highest values in the 2000-year record in the period after 1950. On the basis of this reconstruction and recent instrumental data, Kaufman et al. (2009) concluded that the decade 1999 to 2008 had the warmest summers in the 2000-year record (Kaufman et al., 2009).

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7.2.2.

Arctic ice cores: Holocene climate records from glacierized regions

Ice cores from the larger Arctic ice caps yield records of past environmental conditions in glacierized regions that span the full Holocene. The longest records are from the Canadian Arctic, Mount Logan (Yukon), and Severnaya Zemlya (Koerner and Fisher, 1990, 2002; Kotlyakov et al., 1991; Fisher et al., 1995, 2008; Koerner, 1997; Zdanowicz et al., 2002; Kinnard et al., 2008). The melt layer record from the Agassiz Ice Cap and the oxygen isotope record from Severnaya Zemlya show a temperature maximum in the early Holocene, before 8000 bp, while oxygen isotope records from Arctic Canada typically show a later thermal maximum – as late as 5000 bp on Devon Island (Paterson et al., 1977). Melt layer records show a persistent cooling trend from about 8000 bp until the mid-19th century, after which there is renewed warming. Oxygen isotope records

also show cooling, but the timing of this is delayed until after about 5000 bp in some records (Koerner and Fisher, 2002).

The melt records suggest maximum summer cooling from the Holocene Thermal Maximum of 2.0 to 2.5 °C, while the oxygen isotope records suggest cooling of 1.3 to 3.5 °C (Koerner and Fisher, 2002). Part of the early Holocene warming recorded in the ice core records must be due to the effects of surface lowering due to ice cap thinning, although this would have been partly offset by cooling induced by isostatic uplift of the land surface as the ice cover thinned and shrank (Koerner and Fisher, 2002; Vinther et al., 2009).

Oxygen isotope records from ice cores from smaller ice caps in Arctic Canada, northern Greenland, Svalbard, and the Russian Arctic islands show temporal variability but not the long-term cooling trend apparent in the records from the larger ice caps. This suggests that the small ice caps began to re-grow in the latter part of the Holocene after melting away during the

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GRIP Penny Renland Devon Akademii

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Summer melt, g/y Fraction of years with melt δ18O (0/00)

Figure 7.6. Holocene sections of ice cores from Greenland (GRIP, GISP2, Renland), Arctic Canada (Agassiz, Devon, Penny), and Severnaya Zemlya (Akademii Nauk) showing records of 18O (negative values) and summer melt (positive values), which are proxies for air temperature over the ice caps / ice sheet. Shaded sections suggest conditions warmer than those of today. Source: Koerner and Fisher (2002).

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Thermal Maximum (Hammer et al., 2001; Landvik et al., 2001;

Koerner and Fisher, 2002). The record from the upper 125 m of the ice core from Hans Tausen Iskappe in northern Greenland suggests a warm period between 900 and 1100 ad, and that the period 1700 to 1900 was the coldest in the past 2000 years. There was strong warming from the 1920s to the early 1960s (making the 20th century the warmest part of the record), but no clear trend in temperatures thereafter (Hammer et al., 2001). New oxygen isotope records from Austfonna and Lomonosovfonna in Svalbard span the past 600 to 800 years and show gradual cooling to a minimum in the 19th century, followed by rapid warming around 1900 that made the 20th century the warmest century in the past 600 years (Isaksson et al., 2003, 2005). A 275-year oxygen isotope record from Akademii Nauk Ice Cap on Severnaya Zemlya shows the coldest conditions around 1790, and then continuous warming until 1935. It also shows that the 20th century was the warmest period in this region (Fritzsche et al., 2005). A reconstruction of snow accumulation rates on Lomonosovfonna shows an increase of 25% in the late 1940s (Pohjola et al., 2002). A similar increase is recorded on Severnaya Zemlya after 1935 (Opel et al., 2009).

The oxygen isotope record from Mount Logan is unusual in that it appears to be a proxy for changes in the moisture source region for precipitation, rather than air temperature (Fisher et al., 2008). Major changes in moisture source region seem to be associated with switches between strong, frequent El Niño conditions (associated with meridional patterns of atmospheric flow and water vapor transport towards the Yukon) and strong, frequent La Niña conditions (associated with more zonal flow and moisture transport towards the Pacific Northwest). La Niña periods are associated with reduced precipitation (and snow accumulation) in the southern Yukon, while the reverse is true for El Niño periods. Periods of enhanced meridional flow seem to have occurred around 4200 bp and between 8000 and 7000 bp, while the modern El Niño Southern Oscillation regime seems to have been initiated after 4200 bp (Fisher et al., 2008).

Chemically-based melt indices from the Lomonosovfonna ice core show very high melt in the 12th century (Grinsted et al., 2006), as do melt layer records from the Canadian Arctic.

This suggests a warm episode in medieval times in these parts of the Arctic. This warm episode was followed by a long period of reduced melt lasting to the mid- to late 19th century, after which melt increased sharply. For example, the melt layer record from Akademii Nauk Ice Cap shows maxima in the 1840s, 1880s, and from 1900 to 1970, after which melt layer content dropped sharply (Opel et al., 2009). However, the increase in melt was delayed until about 1925 in the new record from Prince of Wales Icefield, Ellesmere Island (Kinnard et al., 2008).

The first principal component of seven Arctic melt layer records spanning the period 1551 to 1956 explains 34% of the variance in these records. It shows a strong increase in melt layers starting around 1830 and peaking in the mid-20th century (Kinnard et al., 2008) and has a similar form to a multi-proxy summer temperature reconstruction for the Arctic (Overpeck et al., 1997). Melt layer records from multiple sites on the Devon Island Ice Cap show an increase in ice fraction of over 50% after 1989 relative to the period 1963 to 1988 (Colgan and Sharp, 2008), consistent with a shift to increasingly negative surface mass balance after 1987 (Gardner and Sharp, 2007).

7.2.3.

Climate records for glacierized regions of the Arctic derived from instrumental observations and climate re-analysis

Observational records of past surface air temperature from around 70 stations in the Arctic (north of 62° N) since 1875 show large multi-decadal fluctuations. Maxima in mean annual air temperatures occurred in the 1930s and 1940s and in the past two decades, with minima before 1920 and from about 1955 to 1985 (Polyakov et al., 2003). The same trends are apparent for the region 68° to 90° N in the 250-km smoothed air temperature dataset from the Goddard Institute for Space Studies. Mean summer 2-m air temperature anomalies in this dataset show a similar history to mean annual anomalies, but the range of summer anomalies is smaller than for the mean annual air temperature. In addition, there is a period of high mean summer air temperature in the 1950s that is not apparent in the record of mean annual air temperature (Figure 7.7).

Correlations between surface air temperature measured at coastal weather stations and the monthly North Atlantic Oscillation index are positive for a region that includes northern Scandinavia, Svalbard, and the Eurasian Arctic islands, and negative for a region that covers most of Greenland and the Canadian Arctic (Polyakov et al., 2003). Since the late 1990s, surface air temperature anomalies over the Arctic Ocean have become increasingly positive in autumn, especially over those regions where there has been a strong decrease in September sea-ice extent (Serreze et al., 2009). This trend may have resulted in longer melt seasons over glacierized regions of the Canadian and Eurasian Arctic.

In most regions of the Arctic, low winter precipitation means that interannual variability in annual glacier mass balance arises from variability in the summer balance, which is strongly controlled by summer air temperature. Gardner et al. (2009) found positive correlations between mean lower- troposphere summer air temperature from climate re-analyses and surface air temperature measured at the summit elevations

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Figure 7.7. Five-year running means for mean annual and mean summer (June – August) 2-m air temperature anomalies (relative to the mean for 1951 to 1980) for the region 68° to 90° N from the Goddard Institute for Space Studies 250-km smoothed dataset.

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of Arctic ice caps. Sharp and Wang (2009) reported similar correlations between lower tropospheric summer mean air temperature from climate re-analyses and melt season duration on Eurasian Arctic ice caps. For the period 1948 to 2008, NCEP/NCAR (U.S. National Weather Service National Centers for Environmental Prediction and the National Center for Atmospheric Research) Reanalysis (Kalnay et al., 1996) decadal mean 700 hPa air temperatures over glacierized regions of the Arctic in June to August were warmest in either the 1950s (Russian Arctic, northern Ellesmere and Axel Heiberg Islands) or the 2000s (Alaska, Iceland, Svalbard, and the rest of the Canadian Arctic). They were coldest in the 1990s in the Eurasian Arctic and the 1960s and 1970s over Iceland, Alaska, and the Canadian Arctic.

In the more maritime regions of the Arctic (such as Alaska, Iceland, Svalbard, and western Scandinavia), high variability in winter accumulation can also induce interannual variability in the annual mass balance. According to the NCEP/

NCAR Reanalysis, decadal mean winter (September to May) precipitation anomalies were most negative in the 1950s over Alaska, the 1960s over Iceland, and the 1980s over Svalbard.

They were most positive in the 2000s over Alaska and the 1950s over Iceland and Svalbard.

High summer precipitation keeps glacier surface albedo relatively high and reduces summer melt, while low summer precipitation has the opposite effect. In the Eurasian Arctic islands, decadal mean summer precipitation anomalies were typically positive from the 1950s through the 1970s and negative from the 1980s to 2000s. The opposite trend was observed in Iceland.

7.3.

Changes in glacier extent, volume and total mass

• Following the early Holocene warm period, the cooler climate between 6000 years bp and the end of the Little Ice Age resulted in glacier growth and advances across the Arctic.

• In some cases, maximum glacier extents were reached as early as the mid-17th century, but in most regions glaciers were at or close to their maximum extents in the late 19th and/or early 20th centuries.

• Across the Arctic, glaciers began retreating and losing mass in the early 20th century. This trend continued through the mid-20th century, although glaciers in some regions experienced brief episodes of slower retreat, reduced mass loss, or even mass gain. Overall, there were substantial reductions in glacier area and mass across the entire Arctic over the 20th century.

• Mass loss has accelerated across most regions of the Arctic since the mid-1990s.

• In some regions, the timing and duration of episodes of relatively positive or negative mass balance were tied to major changes in atmospheric circulation related to phenomena such as the North Atlantic Oscillation (Iceland, western Scandinavia), the Pacific Decadal Oscillation (western Canada, southern Alaska), and the location and intensity of the summer circumpolar vortex (Arctic Canada).

7.3.1.

Arctic glacier changes during the Holocene

Reconstructions of dated lateral and terminal moraine positions (and associated stratigraphy) are the primary source of information on glacier changes during the Holocene. In western North America, the Cordilleran ice sheet reached its maximum extent and thickness by 16 000 bp. Between 15 000 and 11 000 bp, climate conditions allowed some lobes of the ice sheet to advance (Menounos et al., 2009), and alpine glaciers also advanced in the southern and north-central ranges of the Cordillera, but by 11 000 bp glacier extent was comparable to that at present (Clague et al., 1982). Glacier extent reached a minimum between about 11 500 to 9000 bp, with temperatures 2 to 3 °C above present and generally drier conditions (Kaufman et al., 2004). Subsequently, periods of glacier advance occurred from 8600 to 8200 bp, 7400 to 6500 bp, 4400 to 4000 bp, 3500 to 2800 bp, and 1700 to 1300 bp (Menounos et al., 2009), and maximum Holocene extents were reached during the early 18th or mid-19th centuries ad.

A number of ice caps in Arctic Canada (Devon, Agassiz, Barnes, Penny) contain ice dating from the last glaciation (Paterson et al., 1977; Hooke and Clausen, 1982; Fisher et al., 1983, 1995, 1998; Zdanowicz et al., 2002) and must therefore have survived throughout the Holocene. During deglaciation, the margins of the last Innuitian ice sheet covering the Queen Elizabeth Islands reached the current margins of ice caps on Devon, Ellesmere, and Axel Heiberg Islands between 10 000 and 7500 14C years bp (England et al., 2006). Many of the smaller ice caps and glaciers in this region disappeared during this period (Koerner and Fisher, 2002). Glaciers began to grow again after 1000 14C years bp (Blake, 1981), reaching maximum extents in the late 19th or early 20th centuries. On Baffin Island, the Laurentide Ice Sheet retreated progressively throughout the Holocene to its current margin, that of the Barnes Ice Cap.

Deglaciation was especially rapid around 7000 bp, when the ice cover in Foxe Basin collapsed (Briner et al., 2009). The early Holocene ice extent in this region is poorly constrained but glaciers began to advance again as early as 6000 bp, reaching extents similar to those of the Little Ice Age (the Holocene maximum for alpine glaciers on Baffin Island) by 3500 to 2500 bp (Briner et al., 2009). Small ice caps have been present in the northern interior of Baffin Island for 1600 to 2800 years, but they were absent for much of the middle Holocene. These ice caps started to expand after 1280 ad, and especially after 1450 (Anderson et al., 2008).

Most glaciers in Greenland were smaller during the early Holocene than at present and some areas that are currently occupied by mountain glaciers and ice caps may have become completely ice free in that period (Hammer et al., 2001).

Neoglacial moraines have been reported (from western and southeastern Greenland), but in most regions historical advances were the most extensive after the early Holocene. In all areas apart from northern Greenland (where many ice margins are stationary), local glaciers are currently receding from their historical maximum extents (Kelly and Lowell, 2009).

Iceland was largely ice free during the early Holocene (8000 to 6000 bp), after which Neoglacial cooling began. Glacier extent increased between 4500 and 4000 bp, and again between 3000 and 2500 bp. The most extensive Holocene ice extent was

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reached during the period 1250 to 1900 ad – probably in the latter part of that period between 1700 and 1850 (Geirsdóttir et al., 2009). However, some steep alpine glaciers reached their maximum extents earlier, probably around 2500 bp, when the present ice caps are also likely to have formed. A general glacier recession that began in the 1890s accelerated after 1930. Retreat slowed by the 1960s as a result of cooler summers beginning in the 1940s, and many steep glaciers were advancing in the 1970s. Renewed climate warming since 1985 has led to a new episode of retreat, and all major non-surging outlet glaciers have been retreating since 1995, at rates of up to 100 m/y. The southern outlets of Vatnajökull eroded down into soft sediments during the pre-1890 advance and have been especially susceptible to rapid retreat since then. The volume of Vatnajökull has decreased by about 300 km3 (10%) since 1890 (Björnsson and Pálsson, 2008).

In northern Scandinavia, retreat of the Scandinavian Ice Sheet during the Late Glacial and early Holocene was punctuated by numerous readvances in the period 11 200 to 8000 bp. Most glaciers disappeared completely at some point in the early to mid-Holocene, and ice cover is likely to have reached a minimum between 6600 and 6300 bp (Nesje, 2009).

Subsequently, there was a period of renewed glacier growth during which there were numerous century- to millennium- long ‘Neoglacial’ events. Glaciers in northern Sweden were probably at their Little Ice Age maximum extent in the 17th and early 18th centuries (Karlén, 1988), while those in Norway reached their maxima in the mid-18th century (Winkler, 2003;

Nesje, 2009), or even later in some parts of northern Norway (Bakke et al., 2005). Most Swedish glaciers reached an extent close to their Holocene maximum at the beginning of the 20th century (Holmlund, 1993). Since then, Scandinavian glaciers have generally retreated, particularly since the 1930s (Nesje, 2009). Increased snow accumulation caused some glaciers to advance in the 1990s (Dowdeswell et al., 1997).

In northeast Svalbard, many glaciers started to retreat after the Last Glacial Maximum by about 9500 bp (Blake, 1989). During this period of deglaciation, variations of the glaciers and ice caps in Svalbard were similar to those of glaciers in Scandinavia. However, the Younger Dryas advance was apparently small due to a very dry climate, and the Little Ice Age maximum extent was larger than the Younger Dryas advance (Mangerud and Landvik, 2007). In the early Holocene, glaciers were much smaller than at present and perhaps even completely absent. A period of glacier growth from 3000 to 2500 bp was followed by a warmer period with smaller glaciers (Svendsen and Mangerud, 1997). Dating of relict vegetation found under the cold-based Longyearbreen indicates that late Holocene glacier growth started about 1100 bp. This period of glacier growth culminated in the Little Ice Age maximum with the greatest Holocene glacier extent. Many ice-cored moraines were formed at this time, which may have been as recently as around 1900 (Humlum et al., 2005). Surging of Svalbard glaciers is common and the maximum extent of many glaciers may be a result of surge dynamics (Hagen et al., 1993). Since the Little Ice Age maximum, the glaciers have generally retreated in response to early 20th century warming (Lefauconnier and Hagen, 1991; Hagen et al., 2003a,b).

In the Russian Arctic, outlet glaciers in northwestern Novaya Zemlya retreated from the outer coast by the earliest

Holocene. Prior to about 9500 bp the terminus of the tidewater Shokal’sky Glacier was over 1 km behind its present margin, allowing incursion of the sea (Zeeberg, 2001). A sediment core from Russkaya Gavan spans a period of 800 years and allows reconstruction of the recent history of Shokal’sky Glacier. A noticeable advance around 1400 ad was followed by a major retreat by about 1600. Increased melting in the 1900s is suggested by an increase in sedimentation rates (Polyak et al., 2004). On Franz Josef Land, glaciers were less extensive than at present throughout the period 10 300 to 4400 14C years bp, but re-expanded to their present margins by 2000 14C years bp (Lubinski et al., 1999). A major advance occurred during the past 1000 years, with maxima around 1200, 1400, and in the mid-17th century. Glaciers were more extensive than now at the start of the 20th century; since then there has been widespread glacier retreat.

7.3.2.

Arctic glacier changes in the 20th and 21st centuries

7.3.2.1.

Methods for determining glacier mass balance

Mass balance is the change in mass of a glacier, quantified as the difference between mass gains (accumulation) and mass losses (ablation), over a specified time period. A summary of the processes that determine the mass balance of a glacier is provided in Appendix 7.1 (see also Box 7.1). Over the 20th and beginning of the 21st centuries, many new methods have been developed for measuring glacier mass balance. Each method provides information at unique spatial and temporal resolutions, and samples specific parameters to characterize glacier evolution over time. Important challenges for global mass balance inventories include the integration of these different datasets and quantification of the errors in each approach.

Long-term records of surface mass balance that form the basis of existing global datasets have been acquired from direct sampling of individual glaciers using glaciological methods (Østrem and Brugman, 1991; Cogley, 2005). In this approach, measurements of snow accumulation are acquired from snow probing or digging snow pits to measure snow depth and density (Figure 7.8). Glacier ablation is measured by determining the

Box 7.1. Units for mass balance measurements

Mass balance measurements have units of mass per unit time (see Appendix 7.1). It is also common to report the rate of change of mass per unit area, for example, kg/m2 per year, which is numerically equivalent to the millimetres of water equivalent (w.e.) per year. In this report, mass balances are usually presented in units of gigatonnes per year (Gt/y) or metres of water equivalent per year (m w.e./y). Exceptions are those studies reporting cumulative mass balances, in which case they are not converted to a rate because it is not always clear what exact time period was used for the measurements, and those studies reporting volume changes, in which case they are not converted to a mass equivalent unless the density of the changing snow and ice volume is clearly stated.

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